PSTE 4223 Methodes sismiques Part I Seismic Refraction

PSTE 4223 Methodes sismiques Part I: Seismic Refraction Anne Obermann 2 x 3 h

Overview �Introduction – historical outline �Chapter 1: Fundamental concepts �Chapter 2: Data acquisition and material �Chapter 3: Data interpretation A: Geophysical Interpretation B : Geological Interpretation

Overview �Introduction – historical outline �Chapter 1: Fundamental concepts �Chapter 2: Data acquisition and material �Chapter 3: Data interpretation

Towards refraction seismology � 1885 all that was known about the Earth structure was a vague idea that the density inside had to be much greater than at the surface within 50 years an incredible amount more had been learned using seismology Breakthrough: Seismometer (late 1800’) Instrumental challenge: how to measure ground motion given that the seismometer sits on the ground? Record very small ground motions on the order of 10 -3 cm for distant earthquakes

Towards refraction seismology � Seismometers were developed to record vertical and horizontal motion. � Precise timing, nowadays done using GPS (Global Positioning System) clocks - so that records can be compared between stations. Data are now recorded digitally and made available on the web.

Towards refraction seismology � In 1889, an earthquake in Japan was recorded successfully on several seismometers in Germany. � Milne discovered that observations showed that the time separations between P and S wave arrivals increased with distance from the earthquake. • Thus, the S-P time could be used to measure the distance to the earthquake.

Towards refraction seismology � Next step: Infer the velocity structure of the Earth as a function of depth from the seismograms that were recorded from many different earthquakes (Inverse Problem). � The simplest approach to the inverse problem treats the earth as flat layers of uniform velocity material. The basic geometry is a layer of thickness z, with velocity v 1, overlying a halfspace with a higher velocity v 2. PROBLEM: DATA QUANTITY DEPENDENT ON LARGE EARTHQUAKES – DIFFERENT SOURCES NEEDED!

Towards refraction seismology

Towards refraction seismology � Set out a line or array of geophones � Input a pulse of energy into the ground � Record the arrival times to interpret the velocity structure

Towards refraction seismology

Seismic methods and scales � Controlled source seismology - allows higher resolution studies (m to 100 s km) - can carry out experiments away from tectonic regions � Global seismology (earthquakes) - provides information on global earth structure and large scale velocity anomalies (100 s to 1000 s km) - difficult to image smaller scale structure, particularly away from earthquake source regions

� Seismic refraction - Used to study large scale crustal layering: thickness and velocity Refraction Seismic methods and scales � Seismic reflection - Difficult to determine accurate velocities and depths Reflection - “Imaging” of subsurface reflectors

Applications

Overview �Introduction – historical outline �Chapter 1: Fundamental concepts - Physical notions - Two-layered model - Special cases �Chapter 2: Data acquisition and material �Chapter 3: Data interpretation

Different waves P (compression) + S (shear) waves Surface waves

Huygens Principle Each point along a material acts like a point source of waves. Waves have circular (spherical) wave fronts, these interact constructively (destructively) and produce the wave fronts that we plot as rays.

Snell’s Law Seismic rays obey Snell’s law The angle of incidence equals the angle of reflection. The angle of transmission is related to the angle of incidence through the velocity ratio. Note: the transmitted energy is refracted WR ON G RE P WH RESE Y? NT -> ATIO N !!

Snell’s law: S wave conversion α 1, β 1 α 2, β 2 A conversion from P to S or vice versa can also occur. Still, the angles are determined by the velocity ratios. p is the ray parameter and is constant along each ray.

Snell’s law: Critical Incidence when α 2 > α 1, e 2 > i =>we can increase i. P until e 2 = 90° α 1 α 2 when e 2=90 °, i=ic the critical angle The critically refracted energy travels along the velocity interface at α 2 continually refracting energy back into the upper medium at an angle ic. Head wave

Wave Propagation according to Huygens Principle WRONG REPRESENTATION !!! ->

Wave Propagation according to Huygens Principle

Wave Propagation according to Huygens Principle

Wave Propagation according to Huygens Principle

Seismic Method comparison

Seismic Method comparison Refraction Reflection Typical targets Near-horizontal density contrasts at depths less than ~100 feet Horizontal to dipping density contrasts, and laterally restricted targets such as cavities or tunnels at depths greater than ~50 feet Required Site Conditions Accessible dimensions None greater than ~5 x the depth of interest; unpaved greatly preferred Vertical Resolution 10 to 20 percent of depth 5 to 10 percent of depth Lateral Resolution ~1/2 the geophone spacing Effective Practical Survey Depth 1/5 to 1/4 the maximum shot-geophone separation >50 feet Relative Costs N 3 N-5 N

Two-layered model

Two-layered model Energy from the source can reach the receiver via different paths Direct wave Reflected wave Head wave or Refracted wave

Time-Distance Diagram (Travel Time curves) Think about: � What would a fast velocity look like on this plot? � Why is the direct ray a straight line? � Why must the direct ray plot start at the origin (0, 0)? � Why is the refracted ray a straight line? � Why does the refracted ray not start at the origin? � Why does the reflected ray start at origin? � Why is the reflected ray asymptotic with the direct ray?

Two-layered model Time (t) 1. Direct wave Energy travelling through the top layer, travel-time The travel-time curve for the direct wave is simply a linear function of the seismic velocity and the shot-point to receiver distance Shot Point Direct Ray Slope=1/v 1 Distance (x) Receiver x v 1

Two-layered model -Energy reflecting off the velocity interface. -As the angles of incidence and reflection are equal, the wave reflects halfway between source and receiver. -The reflected ray arrival time is never a first arrival. Shot Point Time (t) 1. Direct wave 2. Reflected wave Distance (x) Receiver Layer 1 v 1 Layer 2 v 2

The travel time curve can be found by noting that x/2 and h 0 form two sides of a right triangle, so Time (t) 2. Reflected wave Distance (x) This curve is a hyperbola, it can be written as “INTERCEPT TIME” GIVES LAYER THICKNESS For x = 0 the reflected wave goes straight up and down, with a travel time of TR(0) = 2 h 1/v 1. At distances much greater than the layer thickness (x >> h), the travel time for the reflected wave asymptotically approaches that of the direct wave. Shot Point Receiver x h 1 Layer 1 v 1 Layer 2 v 2

Two-layered model -Energy refracting across the interface. -Only arrives after critical distance. - Is first arrival only after cross over distance critical distance Time (t) 1. Direct wave 2. Reflected wave 3. Head wave or Refracted wave cross over distance Distance (x) “CRITICAL DISTANCE” ß NO REFRACTED RAYS ic Layer 1 Layer 2 ic ic ic v 1 v 2

3. Head wave or Refracted wave The travel time can be computed by assuming that the wave travels down to the interface such that it impinges at critical angle, then travels just below the interface with the velocity of the lower medium, and finally leaves the interface at the critical angle and travels upwards to the surface. Reminder Show that: . A h 1 X x 0 D ic ic B C v 1 v 2

3. Head wave or Refracted wave The axis intercept time is found by projecting the travel time curve back to x = 0. The intercept time allows a depth estimation. Critical distance xc: distance beyond which critical incidence first occurs. At the critical distance the direct wave arrives before the head wave. At some point, however, the travel time curves cross, and beyond this point the head wave is the first arrival. The crossover distance, xd, where this occurs, is found by setting TD(x) = TH(x) , which yields: The crossover distance is of interest to determine the length of the refraction line.

Travel-time for refracted waves Time (t) critical distance cross over distance Distance (x)

Reminder: Note on Refraction angle Interesting to notice that the higher the velocity contrast, the smaller the refraction angle. V 1 = 1000 m/s V 2 = 5000 m/s λ = 11 ° V 1 = 1000 m/s V 2 = 2000 m/s λ = 30 ° => We can only analyse cases with an increasing velocity function with depth

Summary � v 1 determined from the slope of the direct arrival (straight line passing through the origin) � v 2 determined from the slope of the head wave (straight line first arrival beyond the critical distance) � Layer thickness h 1 determined from the intercept time of the head wave (already knowing v 1 and v 2) h 1

Multiple-layers For multiple layered models we can apply the same process to determine layer thickness and velocity sequentially from the top layer to the bottom.

Multiple-layers � The layer thicknesses are not as easy to find � Recall… Solve for h 1… Now, plug in h 1 and solve the remaining layers one at a time … BEWARE!!! h 1, h 2, are layer thicknesses, not depth to interfaces. So, depth to bottom of layer 3 /top of layer 4 = h 1 + h 2 + h 3

Multiple-layers General formulation

Overview �Introduction �Chapter 1: Fundamental concepts �Chapter 2: Material and data acquisition �Chapter 3: Data interpretation

Material � Geophones � Recording device (Computer, Seismograph) � Source (hammer, explosives) � Battery � Cables � (Geode)

Material: Geophones need a good connection to the ground to decrease the S/N ratio (can be buried)

Material: Cable, Geode

Material: Energy Source � Sledge hammer (Easy to use, cheap) � Buffalo gun (More energy) � Explosives (Much more energy, licence required) � Drop weight (Need a flat area) � Vibrator (Uncommun use for refraction) � Air gun (For lake / marine prospection) Goal: Produce a good energy with high frequencies, Possible investigation depth 10 -50 m You can add (stack) few shots to improve signal/noise ratio

Data acquisition Number of receivers and spacing between them => will define length of the profile and resolution Number of shots to stack (signal to noise ratio) Position of shots

Geophone Spacing / Resolution � Often near surface layers have very low velocities � E. g. soil, subsoil, weathered top layers of rock � These layers are likely of little interest, but due to low velocities, time spent in them may be significant To correctly interpret data these layers must be detected Find compromise between: Geophone array length needs to be 4 -5 times longer than investigation depth Geophone distance cannot be too large, as thin layer won’t be detected

Geophone Spacing / Resolution • This problem is an example of…?

Overview �Introduction �Chapter 1: Fundamental concepts �Chapter 2: Data acquisition and material �Chapter 3: Data processing and interpretation

Record example Dynamite shot recorded using a 120 -channel recording spread

Record example Example of seismic refraction data acquisition where students are using a 'weight-drop' - a 37 kg ball dropped on hard ground from a height of 3 meter - to image the ground to a depth of 1 km

Record example Time Distances

First Break Picking � This is the most important operation, good picking on good data !!!! � A commun problem is the lack of energy, for far offset geophones

First Break Picking –on good data noise

First Break Picking –on poor data noise ?

Travel-time curve How does the inverse shot look like in an planar layered medium? t distance

Reciprocity of travel-times

Assigning different layers

Control of travel-times

Travel time inversion to find best matching underground model

Complete analysis process

Exercice

Some Problems Dipping interfaces Undulating interfaces There are two cases where a seismic interface will not be revealed by a refraction survey. The low velocity layer The hidden layer

Dipping Interfaces • What if the critically refracted interface is not horizontal? � A dipping interface produces a pattern that looks just like a horizontal interface! � Velocities are called “apparent velocities” � What do we do? In this case, velocity of lower layer is underestimated

Dipping Interfaces • To determine if interfaces are dipping… � Shoot lines forward and reversed � If dip is small (< 5 o) you can take average slope � The intercepts will be different at both ends � Implies different thickness Beware: the calculated thicknesses will be perpendicular to the interface, not vertical

Dipping Interfaces � If you shoot down-dip � Slopes on t-x diagram are too steep � Underestimates velocity � May underestimate layer thickness � Converse is true if you shoot up- dip � In both cases the calculated direct ray velocity is the same. • The intercepts tint will also be different at both ends of survey

Problem 1: Low velocity layer If a layer has a lower velocity than the one above… � There can be no critical refraction - The refracted rays are bent towards the normal � There will be no refracted segment on the t-x diagram for the second layer � The t-x diagram to the right will be interpreted as - Two layers - Depth to layer 3 and thickness of layer 1 will be exaggerated Causes: � Sand below clay � Sedimentary rock below igneous rock � (sometimes) sandstone below limestone How Can you Know?

Problem 2: Hidden layer � Recall that the refracted ray eventually overtakes the direct ray (cross over distance). � The second refracted ray may overtake the direct ray first if: � The second layer is thin � The third layer has a much faster velocity

Undulating Interfaces � Undulating interfaces produce non-linear t-x diagrams � There are techniques that can deal with this � delay times & plus minus method � We will see them later…

Detecting Offsets � Offsets are detected as discontinuities in the t-x diagram � Offset because the interface is deeper and D’E’ receives no refracted rays.

Question: To which type of underground model correspond the following traveltime curves? t t distance

Further information http: //www. geomatrix. co. uk/training-videosseismic. php